Chemical buffering of anthropogenic CO2 is the quantitatively most important oceanic process acting as a carbon sink. Carbon dioxide entering the ocean is buffered due to scavenging by the CO32– ions and conversion to HCO3–, that is, the resulting increase in gaseous seawater CO2 concentration is smaller than the amount of CO2 added per unit of seawater volume. Carbon dioxide buffering in seawater is quantified by the Revelle factor (‘buffer factor’, Equation (7.3)), relating the fractional change in seawater pCO2 to the fractional change in total DIC after re-equilibration (Revelle and Suess, 1957; Zeebe and Wolf-Gladrow, 2001):

The lower the Revelle factor, the larger the buffer capacity of seawater. Variability of the buffer factor in the ocean depends mainly on changes in pCO2 and the ratio of DIC to total alkalinity. In the present-day ocean, the buffer factor varies between 8 and 13 (Sabine et al., 2004a; Figure 7.11). With respect to atmospheric pCO2 alone, the inorganic carbon system of the ocean reacts in two ways: (1) seawater re-equilibrates, buffering a significant amount of CO2 from the atmosphere depending on the water volume exposed to equilibration; and (2) the Revelle factor increases with pCO2 (positive feedback; Figure 7.11). Both processes are quantitatively important. While the first is generally considered as a system response, the latter is a feedback process.

The ocean will become less alkaline (seawater pH will decrease) due to CO2 uptake from the atmosphere (see Box 7.3). The ocean’s capacity to buffer increasing atmospheric CO2 will decline in the future as ocean surface pCO2 increases (Figure 7.11a). This anticipated change is certain, with potentially severe consequences.

Increased carbon storage in the deep ocean leads to the dissolution of calcareous sediments below their saturation depth (Broecker and Takahashi, 1978; Feely et al., 2004). The feedback of CaCO3 sediment dissolution to atmospheric pCO2 increase is negative and quantitatively significant on a 1 to 100 kyr time scale, where CaCO3 dissolution will account for a 60 to 70% absorption of the anthropogenic CO2 emissions, while the ocean water column will account for 22 to 33% on a time scale of 0.1 to 1 kyr. In addition, the remaining 7 to 8% may be compensated by long-term terrestrial weathering cycles involving silicate carbonates (Archer et al., 1998). Due to the slow CaCO3 buffering mechanism (and the slow silicate weathering), atmospheric pCO2 will approach a new equilibrium asymptotically only after several tens of thousands of years (Archer, 2005; Figure 7.12).

Figure 7.12. Model projections of the neutralization of anthropogenic CO2 for an ocean-only model, a model including dissolution of CaCO3 sediment and a model including weathering of silicate rocks, (top) for a total of 1,000 GtC of anthropogenic CO2 emissions and (bottom) for a total of 5,000 GtC of anthropogenic CO2. Note that the y-axis is different for the two diagrams. Without CaCO3 dissolution from the seafloor, the buffering of anthropogenic CO2 is limited. Even after 100 kyr, the remaining pCO2 is substantially higher than the pre-industrial value. Source: Archer (2005).

Elevated ambient CO2 levels appear to also influence the production rate of POC by marine calcifying planktonic organisms (e.g., Zondervan et al., 2001). This increased carbon fixation under higher CO2 levels was also observed for three diatom (siliceous phytoplankton) species (Riebesell et al., 1993). It is critical to know whether these increased carbon fixation rates translate into increased export production rates (i.e., removal of carbon to greater depths). Studies of the nutrient to carbon ratio in marine phytoplankton have not yet shown any significant changes related to CO2 concentration of the nutrient utilisation efficiency (expressed through the ‘Redfield ratio’ – carbon:nitrogen:phosphorus:silicon) in organic tissue (Burkhardt et al., 1999).